Hostname: page-component-74656f9884-6grmk Total loading time: 0 Render date: 2026-11-06T09:23:06.722Z Has data issue: false hasContentIssue false
  • English
  • Français

Petrogenesis and geochemical characteristics of Plio-Quaternary alkali basalts from the Qorveh–Bijar volcanic belt, Kurdistan Province, NW Iran

Published online by Cambridge University Press:  17 February 2023

Nafiseh Salehi ,
Ashraf Torkian Open the ORCID record for Ashraf Torkian [Opens in a new window] ,
Tanya Furman  and
Petrus le Roux
Nafiseh Salehi
Affiliation:
Department of Geology, Faculty of Science, Bu-Ali Sina University, Hamedan, Iran
Ashraf Torkian*
Affiliation:
Department of Geology, Faculty of Science, Bu-Ali Sina University, Hamedan, Iran
Tanya Furman
Affiliation:
Department of Geosciences, Pennsylvania State University, University Park, PA, USA
Petrus le Roux
Affiliation:
Department of Geological Sciences, University of Cape Town, Rondebosch 7701, South Africa
*
Author for correspondence: Ashraf Torkian, Email: a-torkian@basu.ac.ir
Rights & Permissions [Opens in a new window]

Abstract

The Pliocene–Quaternary volcanic rocks which outcrop between Qorveh and Bijar are part of post-collisional within-plate volcanic activity in northern Iran. These mafic alkaline rocks form part of the northern arm of the Sanandaj–Sirjan (Hamedan–Tabriz) zone. Thermobarometry on equilibrium clinopyroxene – whole-rock pairs yields pressures and temperatures of 4–6 (±1.8) kbar and 1182–1213 (±27) °C, respectively; olivine – whole-rock (melt) equilibrium thermometry yields crystallization temperatures of 1212–1264 (±27) °C. Field relationships, including the presence of pyroxenitic xenoliths, and geochemical evidence (e.g. high FeO/MnO, and low CaO compared to lavas derived from peridotite sources) suggest a pyroxenitic mantle source for the studied rocks. Variation of trace elements and isotopic ratios (i.e. Ce/Pb, Ba/La, 87Sr/86Sr) indicate that this pyroxenite mantle source was generated by interaction between melted sediments of the subducted Neo-Tethys slab with ambient peridotitic lithospheric mantle. The resulting metasomatized lithosphere is denser and has a lower viscosity than the peridotitic mantle, and tectonic disturbance can cause it to fall into the depths of the mantle. The descending volatile-rich material starts to melt with increasing temperature. Modelling of rare earth element (REE) abundances suggests that <1 % partial melting of the descending pyroxenite could create the Plio-Quaternary alkali basaltic magma of the Qorveh–Bijar. The geochemical evidence for lithospheric foundering, and hence drip magmatism, in the Qorveh–Bijar volcanic belt is supported by seismographic studies indicating thinned lithosphere beneath the study area.

Keywords

partial meltingpyroxenitedrip magmatismisotopic datathermobarometermetasomatized lithosphere

Information

Type
Original Article
Information
Geological Magazine , Volume 160 , Issue 5 , May 2023 , pp. 888 - 904
Copyright
© The Author(s), 2023. Published by Cambridge University Press

1. Introduction

Alkali basaltic magmas are observed in several tectonic environments, including intra-continental, intra-oceanic, post-collisional and arc settings (e.g. Rostami-Hossouri et al. Reference Rostami-Hossouri, Ghasemi, Pang, Shellnutt, Rezaei-Kahkhaei, Miao, Mobasheri, Iizuka, Lee and Lin2020). Kogarko (Reference Kogarko2006) has noted that generally alkaline magmatism is most typical of stable regions where it is controlled by rift structures and occurs along zones marked by abruptly decreasing thickness of the continental lithosphere. These compositions have received considerable attention regarding their geochemistry, petrology and geodynamic evolution (Cebria et al. Reference Cebria, Lopez, Doblas, Oyarzun, Hertogen and Benito2000; Temel et al. Reference Temel, Gourgaud, Alıcı and Bellon2000, Reference Temel, Yürür, Alıcı, Varol, Gourgaud, Bellon and Demirbağ2010; Thompson et al. Reference Thompson, Ottley, Smith, Pearson, Dickin, Morrison, Let and Gibson2005; Xu et al. Reference Xu, Ma, Frey, Feigenson and Liu2005; Jung et al. Reference Jung, Jung, Hoffer and Berndt2006; Kuritani et al. Reference Kuritani, Yokoyama and Nakamura2008, Reference Kuritani, Kimura, Miyamoto, Wei, Shimano, Maeno, Jin and Taniguchi2009; Pilet et al. Reference Pilet, Baker and Stolper2008; Zeng et al. Reference Zeng, Chen, Hofmann, Jiang and Xu2011; Ducea et al. Reference Ducea, Seclaman, Murray, Jianu and Schoenbohm2013; Pang et al. Reference Pang, Chung, Zarrinkoub, Khatib, Mohammadi, Chiu, Chu, Lee and Lo2013; Torkian et al. Reference Torkian, Salehi and Sieble2016; Rostami-Hossouri et al. Reference Rostami-Hossouri, Ghasemi, Pang, Shellnutt, Rezaei-Kahkhaei, Miao, Mobasheri, Iizuka, Lee and Lin2020; Salehi et al. Reference Salehi, Torkian and Furman2020; Verma & Molaei-Yeganeh, Reference Verma and Molaei-Yeganeh2022). The alkaline basalts parental magmas were produced by relatively small degrees of melting (<5 wt %) of their heterogeneous mantle source (Fitton & Dunlop, Reference Fitton and Dunlop1985) and, as such, alkali basalts may be taken as deep probes of enriched domains in the upper mantle (Farmer et al. Reference Farmer, Glazner and Manley2002). Modern geochemical and isotopic data show that alkali–basaltic magmas are formed by melting of enriched reservoirs within the lithospheric and sub-lithospheric mantle (Thompson et al. Reference Thompson, Ottley, Smith, Pearson, Dickin, Morrison, Let and Gibson2005; Xu et al. Reference Xu, Ma, Frey, Feigenson and Liu2005; Jung et al. Reference Jung, Jung, Hoffer and Berndt2006; Kogarko, Reference Kogarko2006; Kogiso & Hirschmann Reference Kogiso and Hirschmann2006; Sobolev et al. Reference Sobolev, Hofmann, Kuzmin, Yaxley, Arndt, Chung, Danyushevsky, Elliott, Frey, Garcia and Gurenko2007; Pilet et al. Reference Pilet, Baker and Stolper2008; Kuritani et al. Reference Kuritani, Kimura, Miyamoto, Wei, Shimano, Maeno, Jin and Taniguchi2009; Zeng et al. Reference Zeng, Chen, Xu, Jiang and Hofmann2010; Ma et al. Reference Ma, Malpas, Xenophontos and Chan2011).

This study explores the petrogenesis of Plio-Quaternary mafic alkaline volcanic rocks that outcrop between the two cities of Qorveh and Bijar. This volcanic episode has been the subject of extensive research, which sets the stage for addressing the origin, age and tectonic settings of the volcanic rocks erupted during this period. The petrogenesis of these volcanic rocks is broadly related to subduction of Neo-Tethyan oceanic crust and continental collision (Malecootyan et al. Reference Malecootyan, Hagh-Nazar, Ghorbani and Emami2007; Kord, Reference Kord2012). Malecootyan et al. (Reference Malecootyan, Hagh-Nazar, Ghorbani and Emami2007) conclude that crustal contamination occurred during the upward movement of magma to the surface and this process was responsible for the distinct compositional characteristics (enrichment in Pb, Rb and Sr and depletion in Nb and Zr) of the Qorveh–Bijar volcanic rocks. Torkian et al. (Reference Torkian, Salehi and Sieble2016) documented the existence of gneissic xenoliths and quartz and alkali feldspar xenocrysts in the NW Qorveh volcanic rocks as evidence of crustal contamination phenomena that may partly overprint the geochemistry of the mantle source. Several authors suggest they derived from an ocean island basalt (OIB)-like mantle source (Moinevaziri & Amin-Sobhani, Reference Moinevaziri and Amin-Sobhani1988; Razavi & Sayyareh, Reference Razavi and Sayyareh2010). Allen et al. (Reference Allen, Kheirkhah, Neill, Emami and McLeod2013) suggested that the high La/Nb and Zr/Hf of the Qorveh–Bijar volcanic rocks indicates a mantle source which was affected by slab-derived fluids. The high Nb concentration and other geochemical features led Azizi et al. (Reference Azizi, Asahara and Tsuboi2014) to interpret these volcanic rocks as high-Nb basalts generated by partial melting of metasomatized mantle associated with adakitic magma. Recent calculations of the parental melt composition based on olivine-hosted melt inclusions demonstrated a pyroxenite source for Quaternary alkaline (Salehi et al. Reference Salehi, Torkian and Furman2020).

Here we present new interpretations based on whole-rock geochemistry (major elements, trace elements and Sr–Nd isotopes) that constrain the contribution of crustal contamination to the genesis of these rocks, as well as highlighting the possible role of subducted oceanic crust in the geochemistry of the mantle source. Detailed mineral chemistry is used to retrieve the intensive variables of the magmatic system. We integrate the petrological and geochemical information derived for the Qorveh–Bijar volcanic belt to provide additional constraints on the melting conditions of the mantle source beneath the Arabian–Eurasian collision zone.

2. Geological setting

The Cenozoic continental collision between the Iranian and Arabian plateaus is manifest in widespread magmatic and metamorphic features in Iran. The subduction of Neo-Tethyan oceanic lithosphere beneath eastern Turkey and Iran initiated during Early Jurassic or Late Triassic time (Dewey et al. Reference Dewey, Pitman, Ryan and Bonnin1973; Berberian & King, Reference Berberian and King1981; Alavi, Reference Alavi1994; Stampfli & Borel, Reference Stampfli and Borel2002; Hassanzadeh & Wernicke, Reference Hassanzadeh and Wernicke2016; Barber et al. Reference Barber, Stockli, Horton and Koshnaw2018; Tavakoli et al. Reference Tavakoli, Davoudian, Shabanian, Azizi, Neubauer, Asahara and Bernroider2020). Subsequent northward motion of the Arabian plate following final closure of Neo-Tethys occurred during the late Oligocene – early Miocene (e.g. Dewey et al. Reference Dewey, Pitman, Ryan and Bonnin1973; Berberian & King, Reference Berberian and King1981; Alavi, Reference Alavi1994; Mouthereau et al. Reference Mouthereau, Lacombe and Vergés2012; Hassanzadeh & Wernicke, Reference Hassanzadeh and Wernicke2016; Barber et al. Reference Barber, Stockli, Horton and Koshnaw2018; Tavakoli et al. Reference Tavakoli, Davoudian, Shabanian, Azizi, Neubauer, Asahara and Bernroider2020) or the Late Cretaceous – Oligocene (Mohajjel & Fergusson Reference Mohajjel and Fergusson2014). The closure of Neo-Tethys has given rise to the East Anatolian and Iranian plateaus to the north and east, respectively, of the Bitlis–Zagros suture (Fig. 1).

Fig. 1.

Late Cenozoic volcanic centres and active faults in Iran. The study area is included in the red rectangle (the rectangle shows the study area from Allen et al. Reference Allen, Kheirkhah, Neill, Emami and McLeod2013).

There are considerable variations in the style and quantity of magmatism after the Arabia–Eurasia collision. Magmatic rocks ranging in age from Miocene to Quaternary are geographically dispersed, volumetrically modest and chemically varied. The complex continental collision zone in western Iran (Fig. 1) consists of the Zagros fold-and-thrust belt (ZFTB), the Sanandaj–Sirjan zone (SaSZ), and the Urumieh–Dokhtar magmatic belt (UDMB) (Berberian & King, Reference Berberian and King1981; Alavi, Reference Alavi1994; Hassanzadeh & Wernicke, Reference Hassanzadeh and Wernicke2016; Tavakoli et al. Reference Tavakoli, Davoudian, Shabanian, Azizi, Neubauer, Asahara and Bernroider2020). The SaSZ can be divided into distinct northern and southern sections (Eftekharnejad, Reference Eftekharnejad1981; Ghasemi & Talbot, Reference Ghasemi and Talbot2006). The northern section is mainly composed of an old island arc and an active continental margin that collided in the Late Jurassic – Early Cretaceous. The southern section consists entirely of metamorphic basement with evidence of polyphase deformation and metamorphism (Azizi & Asahara, Reference Azizi and Asahara2013). The SaSZ has been intruded by A-, S- and I-type granitoid batholiths emplaced from Jurassic to Oligocene time (e.g. Sepahi & Athari, Reference Sepahi and Athari2006; Mansouri-Esfahani et al. Reference Mansouri-Esfahani, Khalili, Kochhar and Gupta2010; Shahbazi et al. Reference Shahbazi, Siebel, Pourmoafee, Ghorbani, Sepahi, Shang and Abedini2010; Torkian & Furman, Reference Torkian and Furman2015; Yeganeh et al. Reference Yeganeh, Torkian, Christiansen and Sepahi2018).

Between the Main Zagros Thrust (MZT) in the southwest and the Tabriz Fault in the northeast, Azizi & Moinevaziri (Reference Azizi and Moinevaziri2009) proposed a subdivision of SaSz in northwestern Iran that is of Cretaceous and Eocene–Miocene to Quaternary age, trending in a NW–SE direction and including three minor volcanic belts: (1) the Sonqor–Baneh volcanic belt (SBVB), (2) the Hamedan–Tabriz volcanic belt (HTVB) and (3) the Cretaceous volcanic belt (SCVB) (see fig. 3 in Azizi & Moinevaziri, Reference Azizi and Moinevaziri2009). The SCVB consists mainly of mafic to intermediate submarine volcanics of calc-alkaline affinity, and the SBVB is composed of basalt, as well as gabbro to dioritic bodies, with extrusive to sub-volcanic magmatic textures and tholeiitic to alkaline affinity.

The HTVB extends across the Hamedan to Tabriz and consists of Miocene to Plio-Quaternary extrusive rocks. The northern part of this belt has Miocene volcanic rocks with adakitic features (Azizi et al. Reference Azizi, Asahara and Tsuboi2014; Lechmann et al. Reference Lechmann, Burg, Ulmer, Guillong and Faridi2018; Torkian et al. Reference Torkian, Furman, Salehi and Veloski2019; Shahbazi et al. Reference Shahbazi, Maghami, Azizi, Asahara, Siebel, Maanijou and Rezai2021). The southern part consists of two different volcanic suites: felsic to intermediate rocks of Miocene age and Plio-Quaternary basalts (Şengör & Kidd, Reference Şengör and Kidd1979; Kheirkhah & Mirnejad, Reference Kheirkhah and Mirnejad2014). Here we investigate mafic volcanic rocks in the HTVB located between the cities of Qorveh and Bijar (i.e. 35° 18′–35° 30′ N, 47° 46′– 47° 59′ E; Fig. 2). The results of K–Ar whole-rock dating in Qorveh–Bijar conducted by Boccaletti et al. (Reference Boccaletti, Innocenti, Manetti, Mazzuoli, Motamed, Pasquare, Radicati di Brozolo and Amin Sobhani1976) suggest that the volcanic activity occurred during the Quaternary, from 1.3 ± 0.08 to 0.5 ± 0.15 Ma.

Fig. 2.

Schematic geological map showing the location of lava flows in the study area. After Emami et al. (Reference Emami, Sadeghi and Omrani1993). The sampling sites are shown. Abbreviations in the diagram are Gh (Ghezelche-Kand), IL (Illanlu), A (Ahmad-Abad), T (Tahmoures) and G (Ghare-Toreh).

3. Field relationships

The Qorveh–Bijar volcanic products comprise bombs, scoria, lapilli tuffs and lava flows with an individual thickness up to several tens of metres (Fig. 3a, b); we refer to these units collectively as the QBB (Qorveh–Bijar basaltic rocks). Cinder cones represent the youngest phase of magmatism in the region, preserving their geological structures over the lava flows. The lava flows cover the argillaceous limestone of Miocene to Pliocene time. There is no significant deposition postdating the lava flows, and three-dimensional structures are exposed through dissection by an external drainage (Fig. 3c).

Fig. 3.

Overview of the products object of this study. (a) Scoria cones, (b) volcanic bombs, (c) scoria and lavas and (d) gneissic xenoliths.

Felsic gneissic xenoliths are frequently observed in the basaltic rocks and some of these xenoliths are larger than 10 cm (Fig. 3d).

4. Material and methods

4.a. Whole-rock geochemistry

Whole-rock major and trace element contents of the studied samples were determined on glassy pills synthetized with the Pt-loop technique at 1600 °C in a chamber furnace installed at the HP-HT Laboratory of Experimental Volcanology and Geophysics of the Istituto Nazionale di Geofisica e Vulcanologia (INGV; Rome, Italy). The pills were then analysed using an electron probe microanalyser (EPMA) Jeol-JXA8200 with combined energy-dispersive spectrometry – wavelength-dispersive spectrometry (EDS-WDS; five spectrometers with 12 crystals) using 15 kV accelerating voltage and 10 nA electric current. A slightly defocused electron beam with a size of 3 μm was used, with a counting time of 5 s on background and 15 son peak. Sodium and potassium were analysed first to prevent alkali migration effects. The accuracy of the microprobe was measured through the analysis of well-characterized synthetic oxides and mineral standards. Based on counting statistics, analytical uncertainties relative to their reported concentrations indicate that precision was better than 5 % for all cations.

Trace element compositions of whole rocks were measured by laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) conducted at the Institute of Geochemistry and Petrology of ETH Zürich (Switzerland) using a 193 nm ArF Excimer laser from Resonetic coupled to a Thermo Element XR ICP-MS. A spot size of 43 μm was used for mineral analyses and reduced to 20 μm for glass analyses; output energy of the laser beam was typically ∼3.5 J cm−2. NIST612 and NIST610 were adopted as external standards for the data reduction. United States Geological Survey (USGS) reference glass GSD-1G was used as a secondary standard to monitor instrument accuracy. When appropriate, major element concentrations from EPMA analyses were used as internal standards. Long-term laboratory reproducibility of homogeneous glass standards indicates precision significantly better than 5 % for elements whose concentration was much greater (i.e. ≥2×) than the detection limit.

4.b. Isotope analysis

Radiogenic isotopic data were obtained at the Department of Earth Science, University of Cape Town (South Africa). Approximately 50 mg of the powdered rock was dissolved in a 4:1 HF/HNO3 acid mixture in sealed Savillex beakers for 48 h, and then the solution was split for determination of both concentration data (Rb, Sr, Nd and Sm), and Sr and Nd isotope ratios. The Sr and Nd fractions for isotope analyses were isolated employing sequential column chemistry (after Pin et al. Reference Pin, Briot, Bassin and Poitrasson1994; Pin & Zalduegui, Reference Pin and Zalduegui1997; Míková & Denková, Reference Míková and Denková2007). The Sr and Nd isotope data were obtained using a Nu Plasma HR mass spectrometer equipped with a DSN-100 desolating nebulizer. All Sr isotopes were referenced to a value of 0.710255 for the bracketing analyses of NIST SRM987. During the analysis, Sr isotope data were corrected for Rb interferences using the measured signal for 85Rb and the natural 85Rb/87Rb ratio, while instrumental mass fractionation was addressed using the exponential law and the 86Sr/88Sr ratio of 0.1194. The Nd isotope values were normalized to 0.512115 for bracketing analyses of JNdi-1. These data were then corrected for Sm and Ce interferences using the signals measured for 147Sm and 140Ce and natural Sm and Ce isotope abundances, while instrumental mass fractionation was addressed using the exponential law and the 146Nd/144Nd ratio of 0.7219.

5. Results

5.a. Petrography

The studied samples are generally fresh and show porphyritic and microlithic textures (Fig. 4a–c). Phenocrysts and microphenocrysts (35–45 vol. %) are represented primarily by clinopyroxene and olivine; in some cases amphibole and biotite are present as accessory phases (Fig. 4e). The groundmass (<35 vol. %) includes microlites of clinopyroxene, acicular plagioclase and opaque minerals (titanomagnetite), all coexisting with glass (∼20 vol. %). Glomeroporphyritic aggregates of olivine and clinopyroxene are observed in some samples.

Fig. 4.

Representative photomicrographs of the studied volcanic rocks. (a) Microlithic porphyry texture; (b) glomeroporphyry texture; (c) porphyry texture; (d) sieve texture of clinopyroxene; (e) amphibole; (f) skeletal olivine. Ol: Olivine; Cpx: Clinopyroxene; Qtz: Quartz; Bt: Biotite; Pl: Plagioclase.

5.b. Mineral chemistry

All mineral compositional data are provided in supplemental files as Tables S1 and S2 (available online at https://doi.org/10.1017/S0016756823000018). Clinopyroxene up to 2 mm is the most abundant mafic mineral phase in all studied rocks. The crystals are commonly euhedral to subhedral and display normal and oscillatory zoning. Some crystal cores and rims show sieve textures with embayments (Fig. 4d). The absence of reaction rims is considered as an indicator of equilibrium between the crystal and the host magma. The clinopyroxenes (Fig. 5a) are classified as diopside to salite with Wo41.3-49.4, En36.4-47.6, Fs6.6–12, Mg# 0.40–0.89 (Mg# expressed as molar Mg/(Mg + Fe+2) where iron is Fe2+ total). Many crystals are slightly zoned, showing increasing TiO2 and FeO concentrations and decreasing MgO contents towards the rims (Fig. 5b, c).

Fig. 5.

(a) Classification scheme of Morimoto et al. (Reference Morimoto, Fabries, Ferguson, Ginzburg, Ross, Seifert and Zussman1988), showing that the pyroxenes are diopside–salite in composition; (b) backscattered electron image microphotograph of clinopyroxene from Illanlu (IL-C7); and (c) mineral compositional variation from rim to rim.

Olivine is the second most abundant phenocryst phase. Crystals are euhedral to subhedral in shape, showing sporadically skeletal and glomeroporphyritic textures (Fig. 4). Some of the olivine phenocrysts display a dissolving–erosion structure, while in other cases the crystals are broken and replaced with iddingsite along fractures and rims. The forsterite content of olivine is variable (Fig. 6a) and generally decreases from core to rim following the normal growth zoning. The highest forsterite content (Fo82–88) is measured in olivine crystals from the Illanlu area. The CaO content of olivine ranges from 0.16 to −2.9 wt %, which is higher than olivine from mantle xenoliths (CaO <0.1 wt %; Thompson & Gibson, Reference Thompson and Gibson2000).

Fig. 6.

(a) Variations of forsterite content in olivine vs Mg# of whole rock. Olivines in equilibrium with the host lavas plot between the dashed lines; (b) Na2O + K2O (wt %) vs SiO2 (wt %) diagram for the QBB rocks (Le Bas et al. Reference Le Bas, Maitre, Streckeisen and Zanettin1986).

5.c. Whole-rock geochemistry

Representative whole-rock (major and trace element) compositions are given in the supplemental file as Table S3 (available online at https://doi.org/10.1017/S0016756823000018). The studied rocks are identified as basanite and phono-tephrite with alkaline affinity in a plot of total alkalis vs SiO2 (Le Bas et al. Reference Le Bas, Maitre, Streckeisen and Zanettin1986) (Fig. 6b); they are generally sodic with Na2O > 2 + K2O.

QBB rocks contain 45.3–48.0 wt % SiO2, 8.1–10.3 wt % MgO, and their Mg# (Mg# = Mg/(Mg + Fe)) ranges from 65 to 72. Variations in Al2O3, Na2O, K2O and SiO2 vs MgO do not define clear trends, and no systematic variations are found between Sr, Nb, La, Th and MgO. However, MgO contents correlate positively with CaO, Ni and Cr (Fig. 7).

Fig. 7.

Bivariate diagrams of selected major and trace elements against MgO (wt %).

Figure 8 shows the chondrite-normalized rare earth elements (REE) and the primitive-mantle-normalized trace element patterns of the QBB rocks. Similar to other intraplate alkaline basalts (Zou et al. Reference Zou, Zindler, Xu and Qi2000; Wilson & Patterson, Reference Wilson and Patterson2001; Shaw et al. Reference Shaw, Baker, Menzies, Thirlwall and Ibrahim2003; Aydin et al. Reference Aydin, Karsli and Chen2008; Asan & Kurt, Reference Asan and Kurt2011; Pang et al. Reference Pang, Chung, Zarrinkoub, Khatib, Mohammadi, Chiu, Chu, Lee and Lo2013), all the samples are enriched in light REE (LREE), exhibiting steep REE patterns (Fig. 8a) with (La/Yb)N values ranging from 33.1 to 68.3. The sub-parallel and tight REE patterns suggest that these volcanic rocks originated from a common mantle source. The QBB are enriched in large-ion lithophile elements (LILE) (Cs: 1.1–3.6 ppm; Sr: 1586–3080 ppm; Pb: 11.2–28.1 ppm), and display negative Nb–Ta anomalies on primitive-mantle normalized abundance diagrams, which is a known characteristic of lavas derived from a mantle source with subduction-modified material or crustal contamination (Fig. 8b).

Fig. 8.

(a) Chondrite-normalized REE diagram and (b) primitive-mantle-normalized trace element diagram for QBB rocks. Normalization values from Sun & McDonough (Reference Sun and McDonough1989).

Whole-rock Nd–Sr isotopic analyses for QBB are reported in Table 1. Initial 87Sr/86Sr and 143 Nd/144Nd ratios of QBB range from 0.70453 to 0.70535 and from 0.512643 to 0.512722 (ϵ Nd +0.23 to +1.76), respectively. The QBB rocks plot close to the composition of the bulk silicate earth and have lower values of 87Sr/86Sr in comparison to the gneissic xenoliths (Azizi et al. Reference Azizi, Asahara and Tsuboi2014) which are considered as continental crust components in the study area (Fig. 9a).

Table 1.

Rb–Sr and Sm–Nd isotopic data for the Qorveh–Bijar basaltic rocks

* The age correction is based on the ages calculated by Boccaletti et al. (Reference Boccaletti, Innocenti, Manetti, Mazzuoli, Motamed, Pasquare, Radicati di Brozolo and Amin Sobhani1976).

Fig. 9.

(a) 87Sr/86Sr vs 143Nd/144Nd. The compositions of Enriched Mantle 1 (EM1) and Enriched Mantle 2 (EM2) reservoirs come from Zindler & Hart (Reference Zindler and Hart1986). The fields for OIBs and high-μ (HIMU) come from Hart (Reference Hart1988), and the composition of gneissic xenoliths comes from Azizi et al. (Reference Azizi, Asahara and Tsuboi2014). (b) The Mg# of clinopyroxene is plotted against the value of cpx-meltKdFe–Mg (Putirka et al. Reference Putirka, Ryerson and Mikaelian2003; Putirka, Reference Putirka2008). Values of cpx-meltKdFe–Mg closely match both the equilibrium ranges of 0.27 ± 0.03 and 0.28 ± 0.08 indicated by Putirka et al. (Reference Putirka, Ryerson and Mikaelian2003) and Putirka (Reference Putirka2008), respectively.

5.d. Intensive parameters

The pressure and temperature conditions of magmas were estimated using the clinopyroxene-melt based thermobarometric models of Putirka et al. (Reference Putirka, Ryerson and Mikaelian2003) and Putirka (Reference Putirka2008), using as input data the compositions of the early-formed crystal cores and the whole-rock analyses (i.e. the original magma compositions). To ascertain whether the clinopyroxene-melt pairs were effectively in equilibrium at the time of crystallization, we employed the equilibrium test of Putirka (Reference Putirka2008) based on Fe–Mg exchange between clinopyroxene core and whole rock (Fig. 9b). As seen in Figure 9b, values of cpx-meltKdFe–Mg closely match, with both the equilibrium ranges of 0.27 ± 0.03 and 0.28 ± 0.08 indicated by Putirka et al. (Reference Putirka, Ryerson and Mikaelian2003) and Putirka (Reference Putirka2008) (their eqs. 32a and 33), respectively. Calculations based on equilibrium clinopyroxene–melt pairs yield pressures and temperatures of 4–6 (±1.8) kbar and 1182–1213 (±27) °C, respectively (Table 2).

Table 2.

Results from clinopyroxene-based thermobarometery, olivine-based thermometry for QBB rocks

Olivine–melt equilibria are particularly useful for liquidus temperature estimates because the Fe–Mg exchange reaction is nearly constant over a wide range of temperature, bulk composition and oxygen fugacity (i.e. Ol-meltKdFe–Mg = 0.30 ± 0.03), and because the olivine Fo content is highly sensitive to the thermal path of magma (e.g. Roeder & Emslie Reference Roeder and Emslie1970; Kuritani et al. Reference Kuritani, Xia, Kimura, Liu, Shimizu, Ushikubo, Zhao, Nakagawa and Yoshimura2019; Rollinson Reference Rollinson2019). Using the olivine-based thermometer approach of Putirka et al. (Reference Putirka, Perfit, Ryerson and Jackson2007) (their eq. 4), we find that Fo85–87 olivine is in equilibrium with the whole-rock data (Fig. 6a), yielding crystallization onset temperatures of 1212–1264 (±27) °C.

6. Discussion

6.a. Fractional crystallization and crustal contamination

Post-melting processes including fractional crystallization and crustal contamination present challenges to deciphering trace element data to determine the nature and composition of the melt source region. We consider the QBB rocks with MgO >10 wt %, Ni ∼300 ppm and Cr >400 ppm to be primary mantle melts. Ni, Cr and CaO contents decrease with decreasing MgO (Fig. 7f–g), consistent with minor fractionation of olivine, clinopyroxene and probably chromian spinel from parental magma. This interpretation is also supported by petrological observations.

Values of Eu/Eu* (0.9–1) and the lack of negative Eu anomalies in chondrite-normalized REE diagrams (Fig. 8) suggest that there is no significant plagioclase fractionation involved in the petrogenesis of the QBB rocks. Many QBB rocks have Ba and Sr abundances that record incompatible behaviour of these elements, consistent with olivine and clinopyroxene fractionation in the absence of plagioclase formation. Following Pang et al. (Reference Pang, Chung, Zarrinkoub, Mohammadi, Yang, Chu, Lee and Lo2012), the absence of negative correlations between Y or Sm, elements with comparatively higher Kd values for amphibole–liquid compared to pyroxene–liquid, and Cr (an index of fractionation) (Fig. 10a and b) indicates that amphibole fractionation was not substantial.

Fig. 10.

(a) Y (ppm) vs Rb (ppm) and (b) Sm vs Rb diagrams for investigating the fractional crystallization for the QBB rocks; (c) Th/Yb vs 87Sr/86Sri and (d) Nb/Th vs ϵ Nd for studied rocks.

Before eruption, intra-plate basalts pass through thick continental crust, creating the possibility that they become contaminated by the crust. The QBB magmas had to pass through the thick continental lithosphere of western Iran (∼110 km; Tunini et al. Reference Tunini, Jimenez-Munt, Fernandez, Verges and Villasenor2014), in which contamination may potentially occur. Indeed, the presence of abundant gneissic xenoliths provides evidence for this process. Numerous mantle xenoliths and xenocrysts are found in the study area; most of the xenoliths are fragmented in appearance with angular edges, suggesting that the host magma ascended too rapidly for them to melt, and thus too rapidly for crustal contamination to play a significant role in the petrogenesis of the QBB (Torkian et al. Reference Torkian, Salehi and Sieble2016; Salehi et al. Reference Salehi, Torkian and Furman2020). The upper continental crust is characterized by enrichment in LILE, depletion in high-field-strength elements (HFSE), high SiO2 (66.6 wt %; Rudnick et al. Reference Rudnick, Gao, Holland and Turekian2003) and enriched Sr–Nd isotopic compositions (87Sr/86Sr = 0.7130, ϵ Nd −15; Gan et al. Reference Gan, Zhang, Barry, He and Wang2018). Consequently, magmas contaminated by continental material should be characterized by elevated SiO2 and LILE concentrations as well as 87Sr/86Sr ratios, but lower HFSE concentrations and 143Nd/144Nd ratios. We emphasize that the QBB lava geochemistry does not display these key features. Further, the lack of systemic positive correlations in plots of Nb/Th – ϵNd and Th/Yb – 87Sr/86Sr (Yu et al. Reference Yu, Chen, Lan, He, Chen and Song2020) implies negligible crustal contamination (Fig. 10c–d).

We employed FC–AFC–FCA® and mixing model software of Ersoy & Helvaci (Reference Ersoy and Helvaci2010) to investigate more fully the possible occurrence of crustal contamination; the model was constrained by the concentrations of incompatible trace elements Nb, Zr and Y in the mafic lavas and the original partition coefficients set in the model (Fig. 11). Primitive mafic lava GH2 with 45.3 wt % SiO2 and 9.3 wt % MgO is assumed as the starting magma composition for AFC modelling. The composition of the contaminant is that of gneissic xenolith sample EGH6 which contains Nb, Zr and Y 5.79 ppm, 98 ppm and 5.77 ppm, respectively (Kord, Reference Kord2012). The investigated ratio of assimilation to fractionation (r) is 0.1, as higher r values would be inconsistent with the primitive MgO contents of the erupted products. Calculated model results plotted in the Zr/Y vs Nb diagram (Fig. 11a) essentially rule out the possibility that the geochemical signature of magmas is due to binary mixing with the continental crust or gneissic xenoliths. Rather, Nb enrichment results from its incompatibility in the fractionating phases of olivine and clinopyroxene. Correlations between 87Sr/86Sr vs Th and Nd143/Nd144 vs Sr further support the trace element modelling, show a limited role for contamination (∼5 %) and make it clear that the enriched nature of the QBB rocks could not result from crustal contamination (Fig. 11 b–c).

Fig. 11.

Geochemical evidence for crustal contamination in Qorveh–Bijar lavas. (a) AFC modelling for the QBB rocks (gneissic xenoliths of Kord, Reference Kord2012); (b) 143Nd/144Nd vs Sr (Moghadam et al. Reference Moghadam, Ghorbani, Khedr, Fazlnia, Chiaradia, Eyuboglu, Santosh, Francisco, Martinez, Gourgaud and Arai2014); (c) 87Sr/86Sr vs Th (ppm) (Ersoy et al. Reference Ersoy, Helvacı, Uysal, Karaoğlu, Palmer and Dindi2012) with modelled assimilation – fractional crystallization pathways (AFC; r = 0.4 curve); (d) Ba/La vs Ce/Pb and (e) Th/Yb vs Ba/La plots to assess the effects of subducted slab materials on the mantle source of QBBs.

The primitive-mantle-normalized incompatible trace element abundance patterns of the QBB are characterized by negative anomalies in Nb and Ta which are a distinctive signature of subduction-related magmas (Wilson, Reference Wilson1989). Therefore, the simple mixing process between a primitive-mantle-derived magma and crustal material cannot be considered a viable mechanism to generate the observed trace element concentrations, and we must consider other factors such as enrichment of the mantle source by subduction components.

6.b. Mantle nature and modelling of melting

Distinguishing the source lithology is pivotal for interpreting the magmatic processes and origin of mantle-derived magmas. This identification can provide important constraints on crustal recycling and/or mantle metasomatism that may have contributed to mantle heterogeneity (Wang et al. Reference Wang, Li, Li, Li, Liu, Long, Zhou and Wang2012, Reference Wang, Li, Li, Pisarevsky and Wingate2014). Peridotites are abundant in the upper mantle, and the vast majority of Earth’s basaltic lavas form through peridotite melting (Hirose & Kushiro, Reference Hirose and Kushiro1993; Walter, Reference Walter1998; Rhodes et al. Reference Rhodes, Huang, Frey, Pringle and Xu2012). Experimental investigations, however, show that partial melts of volatile-free mantle peridotite are unable to match several important geochemical features of intra-plate basalts, including their TiO2, CaO, FeO* and Al2O3 contents (Hirose & Kushiro, Reference Hirose and Kushiro1993; Hirschmann et al. Reference Hirschmann, Kogiso, Baker and Stolper2003; Kogiso et al. Reference Kogiso, Hirschmann and Frost2003). As a result, intra-plate basalts have been suggested to be generated from pyroxenite, peridotite + CO2, and hornblendite source lithologies (Pilet et al. Reference Pilet, Baker and Stolper2008; Ying et al. Reference Ying, Zhang, Tang, Su and Zhou2013). The results of experimental studies illustrate that peridotite and pyroxenite may play a pivotal role in the genesis of basaltic magmas. The existence of pyroxenite in the mantle source of basaltic rocks can be discerned by comparing the major and trace elements chemistry of basaltic magmas with high-pressure experimental products (e.g. Hirschmann et al. Reference Hirschmann, Kogiso, Baker and Stolper2003; Kogiso et al. Reference Kogiso, Hirschmann and Frost2003; Sobolev et al. Reference Sobolev, Hofmann, Kuzmin, Yaxley, Arndt, Chung, Danyushevsky, Elliott, Frey, Garcia and Gurenko2007).

The incompatible element enrichment observed in the QBB rocks could be derived directly from an enriched mantle source (Fig. 8). However, the geodynamic history of the studiy area (SaSZ) leads us to investigate the possible occurrence of mantle metasomatism. It is accepted that slab-derived fluid or melts from ancient Neo-Tethyan oceanic slab subduction beneath the study area could affect the geochemical signature of the mantle source (Agard et al. Reference Agard, Omrani, Jolivet, Whitechurch, Vrielynck, Spakman, Monié, Meyer and Wortel2011). The Ce/Pb ratio is sensitive to the proportion of sediment melt components: subducted sediments incorporated wholesale will increase Ce/Pb values of resulting lavas, whereas fluid components will decrease it because they are rich in fluid-mobile Pb (Tatsumi, Reference Tatsumi2000). We note that in the plot of Ba/La vs Ce/Pb (Fig. 11d), QBB rocks manifested the effects of sediment components in the mantle source. This is an excellent indicator of the type of sedimentary component because sediment-bound Pb is not mobilized by hydrous fluid, whereas it is incompatible during the melting of pelagic sediments (Class et al. Reference Class, Miller, Goldstein and Langmuir2000; Johnson & Plank, Reference Johnson and Plank2000). Moreover, high Th levels are commonly interpreted as reflecting the predominance of a component of subducted pelagic sediments in the magma source (Kirchenbaur et al. Reference Kirchenbaur, Munker and Marchev2009; Kirchenbaur & Munker, Reference Kirchenbaur and Munker2015). The Th/Yb vs Ba/La plot shows that the QBB rocks array supports the involvement of melt components derived from sediments – but not fluids – during the enrichment of the mantle (Fig. 11e).

In addition to incorporating subducted sediments, the mantle source of the QBB experienced metasomatism by silicate melts. Clinopyroxenitic xenoliths have been reported in the Qorveh–Bijar basaltic rocks (Kord, Reference Kord2012) and in Plio-Quaternary alkali basalts of the Marand area in NW Iran (Khezerlou et al. Reference Khezerlou, Amel, Gregoire, Moayyed and Jahangiri2017). These samples provide valuable information on the nature and evolution of the lithospheric mantle in these areas (Downes, Reference Downes1993; Griffin et al. Reference Griffin, Doyle, Ryan, Pearson, Suzanne, Davies, Kivi, Van Achterbergh and Natapov1999; Zhang et al. Reference Zhang, Mahoney, Mo, Ghazi, Milani, Crawford, Guo and Zhao2005; Nasir et al. Reference Nasir, Al-Sayigh, Alharthy and Al-Lazki2006; Ackerman et al. Reference Ackerman, Spacek, Medaris, Hegner, Svojtka and Ulrych2012; Saadat & Stern, Reference Saadat and Stern2012; Ying et al. Reference Ying, Zhang, Tang, Su and Zhou2013).

Complex and diverse mechanisms have been proposed for the formation of pyroxenite veins or zones (Sobolev et al. Reference Sobolev, Hofmann, Kuzmin, Yaxley, Arndt, Chung, Danyushevsky, Elliott, Frey, Garcia and Gurenko2007; Herzberg, Reference Herzberg2011). Mantle pyroxenite can be generated by melting unmodified recycled basaltic crust (stage I pyroxenite) or by the reaction of melted subducted oceanic crust with solid peridotite (stage II pyroxenite; Sobolev et al. Reference Sobolev, Hofmann, Sobolev and Nikogosian2005). As the MgO content of QBB lavas (avg. MgO 9.2 wt %) are expected to be higher from melts of stage I pyroxenite (<8 wt % MgO; Pertermann & Hirschmann, Reference Pertermann and Hirschmann2003), we consider the melting of stage II pyroxenite. The experimental studies of Sobolev et al. (Reference Sobolev, Hofmann, Sobolev and Nikogosian2005) show that eclogite has a lower solidus temperature than peridotite in the lithospheric mantle, therefore eclogite begins melting at higher pressures and greater depth. This melt has high Si concentration and can easily react with olivine-bearing peridotite, converting it to a solid olivine-free pyroxenite. Pyroxenites that result from silicate-melt-modified mantle are often considered the source of oceanic island and intercontinental basaltic rocks (Herzberg, Reference Herzberg2006), and the geochemical characteristics of QBB lavas suggest it is the source of these eruptives.

Mafic melts derived from pyroxenite sources are geochemically distinguishable from melts originating from peridotite sources (Zeng et al. Reference Zeng, Chen, Hofmann, Jiang and Xu2011; Sheldrick et al. Reference Sheldrick, Hahn, Ducea, Stoica, Constenius and Heizler2020); these geochemical signatures are observed consistently in the QBB rocks and suggest contribution from a pyroxenite mantle source. First, melts of pyroxenite have lower CaO contents compared with peridotite-derived basaltic rocks of similar MgO content. While Ca is incompatible with olivine (DCaO l = 0.02; Leeman & Scheidegger, Reference Leeman and Scheidegger1977), the primary constituent of peridotite, it is compatible with clinopyroxene (DCaCpx = 1.8–2.0; Pertermann & Hirschmann, Reference Pertermann and Hirschmann2002). As a result, the CaO content of pyroxenite melts will be lower than that of peridotite melts, as observed in the low CaO content of QBB lavas (Fig. 12a). Second, the QBB lavas have high Fe/Mn values which, following Kogiso & Hirschmann (Reference Kogiso and Hirschmann2001) and Le Roux et al. (Reference Le Roux, Lee and Turner2010), can be attributed to pyroxenite melting (Fig. 12b). Finally, olivine-hosted melt inclusions in QBB mafic lavas manifest higher values of Zn/Fe*10000 than predicted for peridotite-derived melts, supporting a pyroxenite composition for the mantle source of the studied area (Salehi et al. Reference Salehi, Torkian and Furman2020).

Fig. 12.

(a–c) Bulk analyses of QBB mafic rocks display geochemical signatures consistent with melts of a pyroxenite mantle source. (d) Cr (ppm) (a proxy for melting degree) vs normative olivine content of mafic lavas (a proxy for melting pressure) diagram indicates lithospheric drip melting (e.g. Holbig & Grove Reference Holbig and Grove2008; Furman et al. Reference Furman, Nelson and Elkins-Tanton2016). Mafic lava sequences formed through adiabatic upwelling of asthenosphere define a negative correlation. (e) La/Yb vs Sm/Yb melt curves obtained using the non-modal batch melting equation of Shaw (Reference Shaw1970) and mantle sources of both garnet-bearing pyroxenite (Cpx60 + Grt40) and spinel-bearing peridotite (Ol80 + Cpx16 + Grt4) (Kelemen et al. Reference Kelemen, Rilling, Parmentier, Mehl and Hacker2003; Pertermann et al. Reference Pertermann, Hirschmann, Hametner, Günther and Schmidt2004). For both models, the values of C 0 come from the primitive-mantle data reported in Sun and McDonough (Reference Sun and McDonough1989). The QBB rocks plot close to the garnet-bearing pyroxenite-melting curve with ∼1 % of partial melting. Tick marks and numbers along the curves show the degree of partial melting for a given mantle source.

The QBB rocks show LREE enrichments counterbalanced by heavy REE (HREE) depletions; this strong fractionation effect (Fig. 8) suggests that garnet belongs to the phase assemblage of the mantle source (e.g. Coban, Reference Coban2007). Y/Yb values >10 provide an additional clue that garnet is a residual phase in the source region (Ge et al. Reference Ge, Li, Chen and Li2002). We note further that Nb concentrations (>20 ppm) and Nb/Ta values in the studied rocks are high (18–21), consistent with melting in the presence of rutile (Klemme et al. Reference Klemme, Prowatke, Hametner and Günther2005; Liu et al. Reference Liu, Gao, Kelemen and Xu2008).

We explore this question explicitly using Sm/Yb-La/Yb values to distinguish between melts formed in the garnet and spinel stability fields (Fig. 12e). Calculations are consistent with generation of the QBB volcanic rocks by a small degree (about 1 %) of partial melting from a garnet + rutile bearing pyroxenite source (Fig. 12e).

6.c. Geotectonic evolution

Several geological and geophysical studies attribute Iranian and East Anatolian magmatism to the break-off of the southern Neo-Tethyan oceanic slab beneath the Bitlis–Zagros suture and/or delamination of part of the lower lithosphere (e.g. Keskin, Reference Keskin2003; Şengör et al. Reference Şengör, Ozeren, Genc and Zor2003; Molinaro et al. Reference Molinaro, Zeyen and Laurencin2005; Omrani et al. Reference Omrani, Agard, Whitechurch, Benoit, Prouteau and Jolivet2008; Hatzfeld & Molnar Reference Hatzfeld and Molnar2010; Agard et al. Reference Agard, Omrani, Jolivet, Whitechurch, Vrielynck, Spakman, Monié, Meyer and Wortel2011; Chaharlang et al. Reference Chaharlang, Ducea and Ghalamghash2020, Kettanah et al. Reference Kettanah, Abdulrahman, Ismail, MacDonald and Al Humadi2021). Priestley and McKenzie (Reference Priestley and McKenzie2006) suggest that lithosphere thickness in the study area is 150–200 km. Based on this inferred lithosphere thickness, Allen et al. (Reference Allen, Kheirkhah, Neill, Emami and McLeod2013) rejected the process of delamination and suggested that the melting of amphibole- (richterite-)bearing mantle beneath the thickened lithosphere is responsible for the occurrence of melting in this region.

Fichtner et al. (Reference Fichtner, Saygin, Taymaz, Cupillard, Capdeville and Trampert2013) provide a very high-resolution tomographic model (∼10–20 km) at crustal and lithospheric levels which highlights several low-velocity elliptical bodies (∼100–150 km along the shortest axis and 200 km along the longest axis) beneath the study area. These bodies were named ‘compaction pockets’ by Soltanmohammadi et al. (Reference Soltanmohammadi, Grégoire, Rabinowicz, Gerbault, Ceuleneer, Rahgoshay, Bystricky and Benoit2018), who suggested that they could be rising from the mantle transition zone. However, the model suggested by Salehi et al. (Reference Salehi, Torkian and Furman2020) raises the alternative that these bodies could be drips from the lithospheric mantle. Recent investigations by Motavalli-Anbaran et al. (Reference Motavalli-Anbaran, Zeyen, Brunet and Ardestani2011) found that lithospheric thinning (100–120 km) affects the whole of the northern Zagros Mountains including central Iran, relative to a thickness of 180–200 km under the ZFTB and the Persian Gulf. Similarly, Tunini et al. (Reference Tunini, Jimenez-Munt, Fernandez, Verges and Villasenor2014) identified abrupt thinning to c. 140 km under northwestern Iran including the QBB study area. Numerical studies of lithospheric drip and delamination indicate that even this moderate degree of abrupt lithospheric thinning is appropriate for the onset of drip melting. Geochemical modelling of REE abundances in the QBB rocks (Fig. 12d) suggests that the mantle source is garnet ± rutile bearing pyroxenite. Garnet-pyroxenite in the subcontinental lithospheric mantle will be denser than surrounding peridotite and hence gravitationally unstable, so it could delaminate locally and form a metasomatized drip (Elkins-Tanton, Reference Elkins-Tanton2007). As this drip moves downwards it will undergo increased melting as it descends into the hot surrounding asthenosphere. This scenario is in marked contrast to adiabatic upwelling of the asthenosphere, where the degree of melting increases as the depth of melting shallows. The geochemical signature of the studied rocks supports the drip-melting model. Following Holbig & Grove (Reference Holbig and Grove2008), covariation between the amount of normative olivine and Cr concentration in primitive mafic rocks can distinguish between the trends of adiabatic and drip melting (e.g. Furman et al. Reference Furman, Nelson and Elkins-Tanton2016; Gall et al. Reference Gall, Furman, Hanan, Kürkcüoğlu, Sayit, Yürür, Sjoblom, Şen and Şen2021). The QBB rocks follow the trend for increased melting with depth as predicted for drip melting (Fig. 12e).

Edge convection along an abrupt lithospheric boundary can result in the melting of deep lithosphere that is suddenly exposed to heating; this process would develop analogous geochemical signatures in the melts. Undoubtedly, geochemical evidence is not enough to confirm lithospheric drip. However, the oval low-velocity zones could support the notion of partially molten zones within the lithospheric mantle. Their shape suggests they cannot ascend further as they are being compressed rather than rising from the asthenosphere, and we consider these oval-shaped low-velocity zones in the lithosphere to be pieces of foundered lithosphere, i.e., drips. Basaltic melts derived from these drips are likely to be rich in volatiles. They may ascend quickly to the surface along deep-rooted faults, allowing for only a brief stay in magma chambers where they would experience assimilation and fractional crystallization. Among the QBB samples, the low calculated degree of fractionation and the lack of plagioclase in the fractionating assemblage is consistent with this model rather than with a shallow chamber.

7. Conclusion

The Quaternary Qorveh–Bijar basaltic rocks (QBBs) located along a NW–SE trend parallel to the Zagros suture zone are typically alkali basalts with porphyritic, glomeroporphyritic and aphanitic textures. The main crystalline phases are olivine and clinopyroxene. The volcanic rocks show REE and LILE concentrations higher than those of the gneissic xenoliths they carry (which are assumed to represent the continental crust in the study area), indicating that geochemical variations within the QBB suite cannot be attributed to assimilation and/or mixing between primitive magmas and continental crust or gneissic material. High CaO contents, Fe/Mn and Zn/Fe values in the QBB lavas suggest that this mantle source is garnet-bearing pyroxenite in composition. As pyroxenite is denser than peridotitic lithospheric mantle, it is unstable gravitationally and can start to move downwards under its own weight through the process of mantle drip or localized delamination. This model is supported by geophysical data that confirmed the existence of elliptical-shaped low-velocity structures ∼100–200 km in dimension interpreted as melt batches under the study area. Modelling of REE abundances (La/Yb and Sm/Yb) suggests the QBB lavas formed through low degrees of partial melting (∼1 %) from an enriched mantle source in the garnet stability field. This source began to melt during descent in response to increasing temperatures, and the resulting magma ascended along deep-rooted faults, passing through a thick lithosphere where minor assimilation and fractional crystallization took place within the continental crust.

Supplementary material

To view supplementary material for this article, please visit https://doi.org/10.1017/S0016756823000018

Acknowledgements

This research was financially supported by the Bu-Ali Sina University (Iran) {T and S/1395}. The first author expresses her gratitude to the Ministry of Science and Technology of Iran and the Vice-Chancellor for Research and Technology of the Bu-Ali Sina University.

References

Ackerman, L, Spacek, P, Medaris, G Jr, Hegner, E, Svojtka, M and Ulrych, J (2012) Geochemistry and petrology of pyroxenite xenoliths from Cenozoic alkaline basalts, Bohemian Massif. Journal of Geosciences 57, 199–219.Google Scholar
Agard, P, Omrani, J, Jolivet, L, Whitechurch, H, Vrielynck, B, Spakman, W, Monié, P, Meyer, B and Wortel, R (2011) Zagros orogeny: a subduction-dominated process. Geological Magazine 148, 692–725.10.1017/S001675681100046XCrossRefGoogle Scholar
Alavi, M (1994) Tectonics of the Zagros orogenic belt of Iran: new data and interpretations. Tectonophysics 229, 211–38.CrossRefGoogle Scholar
Allen, MB, Kheirkhah, M, Neill, I, Emami, MH and McLeod, CL (2013) Generation of arc and within-plate chemical signatures in collision zone magmatism: Quaternary lavas from Kurdistan Province, Iran. Journal of Petrology 54, 887–911.CrossRefGoogle Scholar
Asan, K and Kurt, H (2011) Petrology and geochemistry of post-collisional early Miocene volcanism in the Karacadaǧ Area (Central Anatolia, Turkey). Acta Geologica Sinica – English Edition 85, 1100–17.10.1111/j.1755-6724.2011.00543.xCrossRefGoogle Scholar
Aydin, F, Karsli, O and Chen, B (2008) Petrogenesis of the Neogene alkaline volcanics with implications for post-collisional lithospheric thinning of the Eastern Pontides, NE Turkey. Lithos 104, 249–66.10.1016/j.lithos.2007.12.010CrossRefGoogle Scholar
Azizi, H and Asahara, Y (2013) Juvenile granite in the Sanandaj–Sirjan Zone, NW Iran: late Jurassic–early Cretaceous arc–continent collision. International Geology Review 55, 1523–40.10.1080/00206814.2013.782959CrossRefGoogle Scholar
Azizi, H, Asahara, Y and Tsuboi, M (2014) Quaternary high-Nb basalts: existence of young oceanic crust under the Sanandaj–Sirjan Zone, NW Iran. International Geology Review 56, 167–86.10.1080/00206814.2013.821268CrossRefGoogle Scholar
Azizi, H and Moinevaziri, H (2009) Review of the tectonic setting of Cretaceous to Quaternary volcanism in northwestern Iran. Journal of Geodynamics 47, 167–79.10.1016/j.jog.2008.12.002CrossRefGoogle Scholar
Barber, DE, Stockli, DF, Horton, BK, Koshnaw, RI (2018) Cenozoic exhumation and foreland basin evolution of the Zagros orogen during the Arabia‐Eurasia collision, western Iran. Tectonics 37, 4396–420.CrossRefGoogle Scholar
Berberian, M and King, GC (1981) Towards a palaeogeography and tectonics evolution of Iran. Canadian Journal of Earth Sciences 18, 210–65.10.1139/e81-019CrossRefGoogle Scholar
Boccaletti, M, Innocenti, F, Manetti, P, Mazzuoli, R, Motamed, A, Pasquare, G, Radicati di Brozolo, F and Amin Sobhani, E (1976) Neogene and Quaternary volcanism of the Bijar Area (Western Iran). Bulletin of Volcanology 40, 122–32.Google Scholar
Cebria, JM, Lopez, RJ, Doblas, M, Oyarzun, R, Hertogen, J and Benito, R (2000) Geochemistry of the Quaternary alkali basalts of Garrotxa (NE Volcanic Province, Spain); a case of double enrichment of the mantle lithosphere. Journal of Volcanology and Geothermal Research 102, 217–35.10.1016/S0377-0273(00)00189-XCrossRefGoogle Scholar
Chaharlang, R, Ducea, MN and Ghalamghash, J (2020) Geochemical evidences for quantifying crustal thickness over time in the Urumieh-Dokhtar magmatic arc (Iran). Lithos 374, 105723.10.1016/j.lithos.2020.105723CrossRefGoogle Scholar
Class, C, Miller, DM, Goldstein, SL and Langmuir, CH (2000) Distinguishing melt and fluid subduction components in Umnak Volcanics, Aleutian Arc. Geochemistry, Geophysics, Geosystems 1, 1–16.10.1029/1999GC000010CrossRefGoogle Scholar
Coban, H (2007) Basalt magma genesis and fractionation in collision-and extension-related provinces: a comparison between eastern, central western Anatolia. Earth-Science Reviews 80, 219–38.10.1016/j.earscirev.2006.08.006CrossRefGoogle Scholar
Dewey, JF, Pitman, WC, Ryan, WB and Bonnin, J (1973) Plate tectonics and the evolution of the Alpine system. Geological Society of America Bulletin 84, 3137–80.2.0.CO;2>CrossRefGoogle Scholar
Downes, H (1993) The nature of the lower continental crust of Europe: petrological and geochemical evidence from xenoliths. Physics of the Earth and Planetary Interiors 79, 195–218.CrossRefGoogle Scholar
Ducea, MN, Seclaman, AC, Murray, KE, Jianu, D and Schoenbohm, LM (2013) Mantle-drip magmatism beneath the Altiplano-Puna plateau, central Andes. Geology 41, 915–18.CrossRefGoogle Scholar
Eftekharnejad, J (1981) Tectonic division of Iran with respect to sedimentary basins. Journal of the Iranian Petroleum Society 82, 19–28 (in Persian).Google Scholar
Elkins-Tanton, LT (2007) Continental magmatism, volatile recycling, and a heterogeneous mantle caused by lithospheric gravitational instabilities. Journal of Geophysical Research 21, 98–112.Google Scholar
Emami, MH, Sadeghi, MMM and Omrani, SJ (1993) Magmatic map of Iran. Map of Iran 1:1,000,000. Tehran: Geological Survey of Iran, internal report.Google Scholar
Ersoy, Y and Helvaci, C (2010) FC–AFC–FCA and mixing modeler: a Microsoft® Excel© spreadsheet program for modelling geochemical differentiation of magma by crystal fractionation, crustal assimilation and mixing. Computers and Geosciences 36, 383–90.10.1016/j.cageo.2009.06.007CrossRefGoogle Scholar
Ersoy, YE, Helvacı, C, Uysal, İ, Karaoğlu, Ö, Palmer, MR and Dindi, F (2012) Petrogenesis of the Miocene volcanism along the İzmir-Balıkesir Transfer Zone in western Anatolia, Turkey: implications for origin and evolution of potassic volcanism in post-collisional areas. Journal of Volcanology and Geothermal Research 241, 21–38.CrossRefGoogle Scholar
Farmer, GL, Glazner, AF and Manley, CR (2002) Did lithospheric delamination trigger late Cenozoic potassic volcanism in the southern Sierra Nevada, California? Geological Society of America Bulletin 114, 754–68.10.1130/0016-7606(2002)114<0754:DLDTLC>2.0.CO;22.0.CO;2>CrossRefGoogle Scholar
Fichtner, A, Saygin, E, Taymaz, T, Cupillard, P, Capdeville, Y and Trampert, J (2013) The deep structure of the North Anatolian fault zone. Earth and Planetary Science Letters 373, 109–17.10.1016/j.epsl.2013.04.027CrossRefGoogle Scholar
Fitton, JG and Dunlop, HM (1985) The cameroon line, West Africa, and its bearing on the origin of oceanic and continental alkali basalt. Earth and Planetary Science Letters 72, 23–38.CrossRefGoogle Scholar
Furman, T, Nelson, WR and Elkins-Tanton, LT (2016) Evolution of the East African rift: drip magmatism, lithospheric thinning and mafic volcanism. Geochimica et Cosmochimica Acta 185, 418–34.CrossRefGoogle Scholar
Gall, H, Furman, T, Hanan, B, Kürkcüoğlu, B, Sayit, K, Yürür, T, Sjoblom, MP, Şen, E and Şen, P (2021) Post-delamination magmatism in south-central Anatolia. Lithos 399, 106299.10.1016/j.lithos.2021.106299CrossRefGoogle Scholar
Gan, CS, Zhang, YZ, Barry, TL, He, JW and Wang, YJ (2018) Jurassic metasomatized lithospheric mantle beneath South China and its implications: geochemical and Sr-Nd isotope evidence from the late Jurassic shoshonitic rocks. Lithos 320, 236–49.CrossRefGoogle Scholar
Ge, X, Li, X, Chen, Z and Li, W (2002) Geochemistry and petrogenesis of Jurassic high Sr/Y low granitoids in eastern China: constraints on crustal thickness. Chinese Science Bulletin 47, 962–80.CrossRefGoogle Scholar
Ghasemi, A and Talbot, CJ (2006) A new tectonic scenario for the Sanandaj–Sirjan Zone (Iran). Journal of Asian Earth Sciences 26, 683–93.10.1016/j.jseaes.2005.01.003CrossRefGoogle Scholar
Griffin, WL, Doyle, BJ, Ryan, CG, Pearson, NJ, Suzanne, YOR, Davies, R, Kivi, K, Van Achterbergh, E and Natapov, LM (1999) Layered mantle lithosphere in the Lac de Gras area, Slave craton: composition, structure and origin. Journal of Petrology 40, 705–27.CrossRefGoogle Scholar
Hart, SR (1988) Heterogeneous mantle domains: signatures, genesis and mixing chronologies. Earth and Planetary Science Letters 90, 273–96.CrossRefGoogle Scholar
Hassanzadeh, J and Wernicke, BP (2016) The Neotethyan Sanandaj-Sirjan zone of Iran as an archetype for passive margin-arc transitions. Tectonics 35, 586–621.CrossRefGoogle Scholar
Hatzfeld, D and Molnar, P (2010) Comparisons of the kinematics and deep structures of the Zagros and Himalaya and of the Iranian and Tibetan plateaus and geodynamic implications. Reviews of Geophysics 48, 1–48.CrossRefGoogle Scholar
Herzberg, C (2006) Petrology and thermal structure of the Hawaiian plume from Mauna Kea volcano. Nature 444, 605–9.CrossRefGoogle ScholarPubMed
Herzberg, C (2011) Identification of source lithology in the Hawaiian and Canary Islands: implications for origins. Journal of Petrology 52, 113–46.CrossRefGoogle Scholar
Hirose, K and Kushiro, I (1993) Partial melting of dry peridotites at high pressures: determination of compositions of melts segregated from peridotite using aggregates of diamond. Earth and Planetary Science Letters 114, 477–89.CrossRefGoogle Scholar
Hirschmann, MM, Kogiso, T, Baker, MB and Stolper, EM (2003) Alkalic magmas generated by partial melting of garnet pyroxenite. Geology 31, 481–4.10.1130/0091-7613(2003)031<0481:AMGBPM>2.0.CO;22.0.CO;2>CrossRefGoogle Scholar
Holbig, ES and Grove, TL (2008) Mantle melting beneath the Tibetan Plateau: experimental constraints on ultra potassic magmatism. Journal of Geophysical Research 113, 48–59.CrossRefGoogle Scholar
Johnson, MC and Plank, T (2000) Dehydration and melting experiments constrain the fate of subducted sediments. Geochemistry, Geophysics, Geosystems 1, 1–28.CrossRefGoogle Scholar
Jung, C, Jung, S, Hoffer, E and Berndt, J (2006) Petrogenesis of Tertiary mafic alkaline magmas in the Hocheifel: Germany. Journal of Petrology 47, 1637–71.CrossRefGoogle Scholar
Kelemen, PB, Rilling, JL, Parmentier, EM, Mehl, L and Hacker, BR (2003) Thermal structure due to solid-state flow in the mantle wedge beneath arcs. Geophysical Monograph – American Geophysical Union 138, 293–311.Google Scholar
Keskin, M (2003) Magma generation by slab steepening and breakoff beneath a subduction-accretion complex: an alternative model for collision-related volcanism in Eastern Anatolia, Turkey. Geophysical Research Letters 30, 1–4.CrossRefGoogle Scholar
Kettanah, YA, Abdulrahman, AS, Ismail, SA, MacDonald, DJ and Al Humadi, H (2021) Petrography, mineralogy, and geochemistry of the Hemrin Basalt, Northern Iraq: implications for petrogenesis and geotectonics. Lithos 390, 106109.CrossRefGoogle Scholar
Kheirkhah, M and Mirnejad, H (2014) Volcanism from an active continental collision zone: a case study on most recent lavas within Turkish-Iranian plateau. Journal of Tethys 2, 81–92.Google Scholar
Khezerlou, AA, Amel, N, Gregoire, M, Moayyed, M and Jahangiri, A (2017) Geochemistry and mineral chemistry of pyroxenite xenoliths and host volcanic alkaline rocks from North West of Marand (NW Iran). Mineralogy and Petrology 111, 865–85.CrossRefGoogle Scholar
Kirchenbaur, A, Munker, C and Marchev, P (2009) The HFSE budget of post-collisional high-K basalts and shoshonites. Geochimica et Cosmochimica Acta 73, 1417–65.Google Scholar
Kirchenbaur, M and Munker, C (2015) The behaviour of the extended HFSE group (Nb, Ta, Zr, Hf, W, Mo) during the petrogenesis of mafic K-rich lavas: the Eastern Mediterranean case. Geochimica et Cosmochimica Acta 165, 178–99.CrossRefGoogle Scholar
Klemme, S, Prowatke, S, Hametner, K and Günther, D (2005) Partitioning of trace elements between rutile and silicate melts: implications for subduction zones. Geochimica et Cosmochimica Acta 69, 2361–71.CrossRefGoogle Scholar
Kogarko, LN (2006) Alkaline magmatism and enriched mantle reservoirs: mechanisms, time, and depth of formation. Geochemistry International 44, 3–10.CrossRefGoogle Scholar
Kogiso, T and Hirschmann, MM (2001) Experimental study of clinopyroxenite partial melting and the origin of ultra-calcic melt inclusions. Contributions to Mineralogy and Petrology 142, 347–60.10.1007/s004100100295CrossRefGoogle Scholar
Kogiso, T and Hirschmann, MM (2006) Partial melting experiments of bimineralic eclogite and the role of recycled mafic oceanic crust in the genesis of ocean island basalts. Earth and Planetary Science Letters 249, 188–99.CrossRefGoogle Scholar
Kogiso, T, Hirschmann, MM and Frost, DJ (2003) High-pressure partial melting of garnet pyroxenite: possible mafic lithologies in the source of ocean island basalts. Earth and Planetary Science Letters 216, 603–17.CrossRefGoogle Scholar
Kord, M (2012) Study of ultramafic and gneissic enclaves in basaltic rocks, NE Qorveh, Kurdistan. MSc thesis. Bu-Ali Sina University, Hamedan, Iran, 137 pp. Published thesis.Google Scholar
Kuritani, T, Kimura, JI, Miyamoto, T, Wei, H, Shimano, T, Maeno, F, Jin, X and Taniguchi, H (2009) Intraplate magmatism related to deceleration of upwelling asthenospheric mantle: implications from the Changbaishan shield basalts, northeast China. Lithos 112, 247–58.CrossRefGoogle Scholar
Kuritani, T, Xia, QK, Kimura, JI, Liu, J, Shimizu, K, Ushikubo, T, Zhao, D, Nakagawa, M and Yoshimura, S (2019) Buoyant hydrous mantle plume from the mantle transition zone. Scientific Reports 9, 1–7.CrossRefGoogle ScholarPubMed
Kuritani, T, Yokoyama, T and Nakamura, E (2008) Generation of rear-arc magmas induced by influx of slab-derived supercritical liquids: implications from alkali basalt lavas from Rishiri Volcano, Kurile arc. Journal of Petrology 49, 1319–42.CrossRefGoogle Scholar
Le Bas, ML, Maitre, RL, Streckeisen, A and Zanettin, B (1986) IUGS subcommission on the systematics of igneous rocks. A chemical classification of volcanic rocks based on the total alkali-silica diagram. Journal of Petrology 27, 745–50.10.1093/petrology/27.3.745CrossRefGoogle Scholar
Le Roux, V, Lee, CT and Turner, SJ (2010) Zn/Fe systematics in mafic and ultramafic systems: implications for detecting major element heterogeneities in the Earth’s mantle. Geochimica et Cosmochimica Acta 74, 2779–96.CrossRefGoogle Scholar
Lechmann, A, Burg, JP, Ulmer, P, Guillong, M and Faridi, M (2018) Metasomatized mantle as the source of Mid-Miocene-Quaternary volcanism in NW-Iranian Azerbaijan: geochronological and geochemical evidence. Lithos 304, 311–28.CrossRefGoogle Scholar
Leeman, WP and Scheidegger, KF (1977) Olivine/liquid distribution coefficients and a test for crystal-liquid equilibrium. Earth and Planetary Science Letters 35, 247–57.CrossRefGoogle Scholar
Liu, Y, Gao, S, Kelemen, PB and Xu, W (2008) Recycled crust controls contrasting source compositions of Mesozoic and Cenozoic basalts in the North China craton. Geochimica et Cosmochimica Acta 72, 2349–76.CrossRefGoogle Scholar
Ma, GS-K, Malpas, J, Xenophontos, C and Chan, GH-N (2011) Petrogenesis of latest Miocene-Quaternary continental intraplate volcanism along the northern Dead Sea Fault System (Al Ghab-Homs volcanic field), western Syria: evidence for lithosphere-asthenosphere interaction. Journal of Petrology 52, 401–30.CrossRefGoogle Scholar
Malecootyan, S, Hagh-Nazar, S, Ghorbani, M and Emami, MH (2007) Magmatic evolution in Quaternary basaltic rocks in Ghorveh–Takab axis. Journal of Geoscience 16, 166–78 (in Persian with English abstract).Google Scholar
Mansouri-Esfahani, MM, Khalili, M, Kochhar, N and Gupta, LN (2010) A-type granite of the Hasan Robat area (NW of Isfahan, Iran) and its tectonic significance. Journal of Asian Earth Sciences 37, 207–18.CrossRefGoogle Scholar
Míková, J and Denková, P (2007) Modified chromatographic separation scheme for Sr and Nd isotope analysis in geological silicate samples. Journal of Geosciences 52, 221–6.Google Scholar
Moghadam, HS, Ghorbani, G, Khedr, MZ, Fazlnia, N, Chiaradia, M, Eyuboglu, Y, Santosh, M, Francisco, CG, Martinez, ML, Gourgaud, A and Arai, S (2014) Late Miocene K-rich volcanism in the Eslamieh Peninsula (Saray), NW Iran: implications for geodynamic evolution of the Turkish–Iranian High Plateau. Gondwana Research 26, 1028–50.CrossRefGoogle Scholar
Mohajjel, M and Fergusson, CL (2014) Jurassic to Cenozoic tectonics of the Zagros Orogen in northwestern Iran. International Geology Review 56, 263–87.CrossRefGoogle Scholar
Moinevaziri, H and Amin-Sobhani, H (1988) Study on Young Volcanoes of Qorveh-Takab Area. Tehran: Tarbiat Moalem University Publications.Google Scholar
Molinaro, M, Zeyen, H and Laurencin, X (2005) Lithospheric structure beneath the south-eastern Zagros Mountains, Iran: recent slab break-off? Terra Nova 17, 1–6.CrossRefGoogle Scholar
Morimoto, N, Fabries, J, Ferguson, AK, Ginzburg, IV, Ross, M, Seifert, FA and Zussman, J (1988) Nomenclature of pyroxenes. Mineralogical Magazine 52, 535–50.10.1180/minmag.1988.052.367.15CrossRefGoogle Scholar
Motavalli-Anbaran, SH, Zeyen, H, Brunet, MF and Ardestani, VE (2011) Crustal and lithospheric structure of the Alborz Mountains, Iran, and surrounding areas from integrated geophysical modelling. Tectonics 30, 1–16.CrossRefGoogle Scholar
Mouthereau, F, Lacombe, O and Vergés, J (2012) Building the Zagros collisional orogen: timing, strain distribution and the dynamics of Arabia/Eurasia plate convergence. Tectonophysics 532, 27–60.CrossRefGoogle Scholar
Nasir, S, Al-Sayigh, A, Alharthy, A and Al-Lazki, A (2006) Geochemistry and petrology of tertiary volcanic rocks and related ultramafic xenoliths from the central and eastern Oman Mountains. Lithos 90, 49–270.CrossRefGoogle Scholar
Omrani, J, Agard, P, Whitechurch, H, Benoit, M, Prouteau, G and Jolivet, L (2008) Arc-magmatism and subduction history beneath the Zagros Mountains, Iran: a new report of adakites and geodynamic consequences. Lithos 106, 380–98.10.1016/j.lithos.2008.09.008CrossRefGoogle Scholar
Pang, KN, Chung, S, Zarrinkoub, MH, Khatib, MM, Mohammadi, SS, Chiu, H, Chu, CH, Lee, H and Lo, C (2013) Eocene–Oligocene post-collisional magmatism in the Lut–Sistan region, eastern Iran: magma genesis and tectonic implications. Lithos 180, 234–51.CrossRefGoogle Scholar
Pang, KN, Chung, SL, Zarrinkoub, MH, Mohammadi, SS, Yang, HM, Chu, CH, Lee, HY and Lo, CH (2012) Age, geochemical characteristics and petrogenesis of Late Cenozoic intraplate alkali basalts in the Lut–Sistan region, eastern Iran. Chemical Geology 306, 40–53.CrossRefGoogle Scholar
Pertermann, M and Hirschmann, MM (2002) Trace-element partitioning between vacancy-rich eclogitic clinopyroxene and silicate melt. American Mineralogist 87, 1365–76.CrossRefGoogle Scholar
Pertermann, M and Hirschmann, MM (2003) Partial melting experiments on a MORB-like pyroxenite between 2 and 3 GPa: Constraints on the presence of pyroxenite in basalt source regions from solidus location and melting rate. Journal of Geophysical Research 108, 1–12.CrossRefGoogle Scholar
Pertermann, M, Hirschmann, MM, Hametner, K, Günther, D and Schmidt, MW (2004) Experimental determination of trace element partitioning between garnet and silica-rich liquid during anhydrous partial melting of MORB-like eclogite. Geochemistry, Geophysics, and Geosystems 5, 2173–2201.CrossRefGoogle Scholar
Pilet, S, Baker, MB and Stolper, EM (2008) Metasomatized lithosphere and the origin of alkaline lavas. Science 320, 916–19.CrossRefGoogle ScholarPubMed
Pin, C, Briot, D, Bassin, C and Poitrasson, F (1994) Concomitant separation of strontium and samarium-neodymium for isotopic analysis in silicate samples, based on specific extraction chromatography. Analytica Chimica Acta 298, 209–17.CrossRefGoogle Scholar
Pin, C and Zalduegui, JS (1997) Sequential separation of light rare-earth elements, thorium and uranium by miniaturized extraction chromatography: application to isotopic analyses of silicate rocks. Analytica Chimica Acta 339, 79–89.CrossRefGoogle Scholar
Priestley, K and McKenzie, D (2006) The thermal structure of the lithosphere from shear wave velocities. Earth and Planetary Science Letters 244, 285–301.CrossRefGoogle Scholar
Putirka, K, Perfit, M, Ryerson, FJ and Jackson, MG (2007) Ambient and excess mantle temperatures, olivine thermometry, and active vs. passive upwelling. Chemical Geology 241, 177–206.10.1016/j.chemgeo.2007.01.014CrossRefGoogle Scholar
Putirka, KD (2008) Thermometers and barometers for volcanic systems. Reviews in Mineralogy and Geochemistry 69, 61–120.CrossRefGoogle Scholar
Putirka, KD, Ryerson, FJ and Mikaelian, H (2003) New igneous thermobarometers for mafic and evolved lava compositions, based on clinopyroxene + liquid equilibria. American Mineralogist 88, 1542–54.CrossRefGoogle Scholar
Razavi, MH and Sayyareh, A (2010) Properties of young volcanic rocks in southeast of Bijar. Journal of Geoscience 19, 151–6.Google Scholar
Rhodes, JM, Huang, S, Frey, FA, Pringle, M and Xu, G (2012) Compositional diversity of Mauna Kea shield lavas recovered by the Hawaii Scientific Drilling Project: inferences on source lithology, magma supply, and the role of multiple volcanoes. Geochemistry, Geophysics, Geosystems 13, 105747.CrossRefGoogle Scholar
Roeder, PL and Emslie, R (1970) Olivine-liquid equilibrium. Contributions to Mineralogy and Petrology 29, 275–89.CrossRefGoogle Scholar
Rollinson, H (2019) Dunites in the mantle section of the Oman ophiolite: the boninite connection. Lithos 334–335, 1–7.CrossRefGoogle Scholar
Rostami-Hossouri, M, Ghasemi, H, Pang, KN, Shellnutt, JG, Rezaei-Kahkhaei, M, Miao, L, Mobasheri, M, Iizuka, Y, Lee, HY and Lin, TH (2020) Geochemistry of continental alkali basalts in the Sabzevar region, northern Iran: implications for the role of pyroxenite in magma genesis. Contributions to Mineralogy and Petrology 175, 1–22.CrossRefGoogle Scholar
Rudnick, RL, Gao, S, Holland, HD and Turekian, KK (2003) Composition of the continental crust. The Crust 3, 1–64.Google Scholar
Saadat, S and Stern, CR (2012) Petrochemistry of a xenolith-bearing Neogene alkali olivine basalt from north-eastern Iran. Journal of Volcanology and Geothermal Research 225, 13–29.CrossRefGoogle Scholar
Salehi, N, Torkian, A and Furman, T (2020) Olivine-hosted melt inclusions in Pliocene–Quaternary lavas from the Qorveh–Bijar volcanic belt, western Iran: implications for source lithology and cooling history. International Geology Review 62, 1828–44.CrossRefGoogle Scholar
Şengör, A, Ozeren, S, Genc, T and Zor, E (2003) East Anatolian high plateau as a mantle-supported, north-south shortened domal structure. Geophysical Research Letters 30, 1–24.CrossRefGoogle Scholar
Şengör, AMC and Kidd, WSF (1979) Post-collisional tectonics of the Turkish-Iranian plateau and a comparison with Tibet. Tectonophysics 55, 361–76.CrossRefGoogle Scholar
Sepahi, AA and Athari, SF (2006) Petrology of major granitic plutons of the northwestern part of the Sanandaj-Sirjan Metamorphic Belt, Zagros Orogen, Iran: with emphasis on A-type granitoids from the SE Saqqez area. Neues Jahrbuch für Mineralogie – Abhandlungen 183, 93–106.CrossRefGoogle Scholar
Shahbazi, H, Maghami, YT, Azizi, H, Asahara, Y, Siebel, W, Maanijou, M and Rezai, A (2021) Zircon U–Pb ages and petrogenesis of late Miocene adakitic rocks from the Sari Gunay gold deposit, NW Iran. Geological Magazine 158, 1733–55.CrossRefGoogle Scholar
Shahbazi, H, Siebel, W, Pourmoafee, M, Ghorbani, M, Sepahi, AA, Shang, CK and Abedini, MV (2010) Geochemistry and U–Pb zircon geochronology of the Alvand plutonic complex in Sanandaj–Sirjan Zone (Iran): new evidence for Jurassic magmatism. Journal of Asian Earth Sciences 39, 668–83.CrossRefGoogle Scholar
Shaw, DM (1970) Trace element fractionation during anataxis. Geochimica et Cosmochimica Acta 34, 237–43.CrossRefGoogle Scholar
Shaw, JE, Baker, JA, Menzies, MA, Thirlwall, MF and Ibrahim, KM (2003) Petrogenesis of the largest intraplate volcanic field on the Arabian Plate (Jordan): a mixed lithosphere–asthenosphere source activated by lithospheric extension. Journal of Petrology 44, 1657–79.CrossRefGoogle Scholar
Sheldrick, TC, Hahn, G, Ducea, MN, Stoica, AM, Constenius, K and Heizler, M (2020) Peridotite versus pyroxenite input in Mongolian Mesozoic-Cenozoic lavas, and dykes. Lithos 376, 105747.CrossRefGoogle Scholar
Sobolev, AV, Hofmann, AW, Kuzmin, DV, Yaxley, GM, Arndt, NT, Chung, SL, Danyushevsky, LV, Elliott, T, Frey, FA, Garcia, MO and Gurenko, AA (2007) The amount of recycled crust in sources of mantle-derived melts. Science 316, 412–7.CrossRefGoogle ScholarPubMed
Sobolev, AV, Hofmann, AW, Sobolev, SV and Nikogosian, IK (2005) An olivine-free mantle source of Hawaiian shield basalts: Nature 434, 590–7.CrossRefGoogle ScholarPubMed
Soltanmohammadi, A, Grégoire, M, Rabinowicz, M, Gerbault, M, Ceuleneer, G, Rahgoshay, M, Bystricky, M and Benoit, M (2018) Transport of volatile-rich melt from the mantle transition zone via compaction pockets: implications for mantle metasomatism and the origin of alkaline lavas in the Turkish–Iranian plate. Journal of Petrology 59, 2273–310.CrossRefGoogle Scholar
Stampfli, GM and Borel, GD (2002) A plate tectonic model for the Paleozoic and Mesozoic constrained by dynamic plate boundaries and restored synthetic oceanic isochrons. Earth and Planetary Science Letters 196, 17–33.CrossRefGoogle Scholar
Sun, SS and McDonough, QF (1989) Chemical and isotopic systematic of oceanic basalts; implications for mantle compositions and processes. In Magmatism in the Ocean Basins (eds AD Saunders and MJ Norry), pp. 312–45. Geological Society of London, Special Publication no. 42,.CrossRefGoogle Scholar
Tatsumi, Y (2000) Slab melting: its role in continental crust formation and mantle evolution. Geophysical Research Letters 27, 3941–4.CrossRefGoogle Scholar
Tavakoli, N, Davoudian, AR, Shabanian, N, Azizi, H, Neubauer, F, Asahara, Y and Bernroider, M (2020) Zircon U-Pb dating, mineralogy and geochemical characteristics of the gabbro and gabbro-diorite bodies, Boein–Miandasht, western Iran. International Geology Review 62, 1658–76.CrossRefGoogle Scholar
Temel, A, Gourgaud, A, Alıcı, P and Bellon, H (2000) The role of asthenospheric mantle in the generation of Tertiary basaltic alkaline volcanism in the Polatlı – Ankara region, central Anatolia, Turkey: constraints from major-element, traceelement and Sr–Nd isotopes, Goldschmidt 2000, September 3–8th. Journal of Conference Abstracts 5, 989.Google Scholar
Temel, A, Yürür, T, Alıcı, P, Varol, E, Gourgaud, A, Bellon, H and Demirbağ, H (2010) Alkaline series related to Early-Middle Miocene intra-continental rifting in a collision zone: an example from Polatlı, Central Anatolia, Turkey. Journal of Asian Earth Sciences 38, 289–306.CrossRefGoogle Scholar
Thompson, RN and Gibson, SA (2000) Transient high temperatures in mantle plume heads inferred from magnesian olivines in Phanerozoic picrites. Nature 407, 502–6.CrossRefGoogle ScholarPubMed
Thompson, RN, Ottley, CJ, Smith, PM, Pearson, DG, Dickin, AP, Morrison, MA, Let, PT and Gibson, SA (2005) Source of the Quaternary alkalic basalts, picrites and basanites of the Potrillo Volcanic Field, New Mexico, USA: lithosphere or convecting mantle? Journal of Petrology 46, 1603–43.CrossRefGoogle Scholar
Torkian, A and Furman, T (2015) The significance of mafic microgranular enclaves in the petrogenesis of the Qorveh Granitoid Complex, northern Sanandaj-Sirjan Zone, Iran. Neues Jahrbuch für Mineralogie – Abhandlungen 192, 117–33.CrossRefGoogle Scholar
Torkian, A, Furman, T, Salehi, N and Veloski, K (2019) Petrogenesis of adakites from the Sheyda volcano, NW Iran. Journal of African Earth Sciences 150, 194–204.CrossRefGoogle Scholar
Torkian, A, Salehi, N and Sieble, W (2016) Geochemistry and petrology of basaltic lavas from NE-Qorveh, Kurdistan province, Western Iran. Journal of Mineral Chemistry 193, 95–112.Google Scholar
Tunini, L, Jimenez-Munt, I, Fernandez, M, Verges, J and Villasenor, A (2014) Lithospheric mantle heterogeneities beneath the Zagros Mountains and the Iranian Plateau: a petrological-geophysical study. Geophysical Journal International 200, 596–614.CrossRefGoogle Scholar
Verma, SP and Molaei-Yeganeh, T (2022) Tectonic settings of the Plio-Quaternary volcanism in Iran from multidimensional and multielement solutions. Geological Journal 57, 410–24.CrossRefGoogle Scholar
Walter, MJ (1998) Melting of garnet peridotite and the origin of komatiite and depleted lithosphere. Journal of Petrology 39, 29–60.CrossRefGoogle Scholar
Wang, XC, Li, ZX, Li, J, Pisarevsky, SA and Wingate, MT (2014) Genesis of the 1.21 Ga Marnda Moorn large igneous province by plume–lithosphere interaction. Precambrian Research 241, 85–103.CrossRefGoogle Scholar
Wang, XC, Li, ZX, Li, XH, Li, J, Liu, Y, Long, WG, Zhou, JB and Wang, F (2012) Temperature, pressure, and composition of the mantle source region of Late Cenozoic basalts in Hainan Island, SE Asia: a consequence of a young thermal mantle plume close to subduction zones?. Journal of Petrology 53, 177–233.CrossRefGoogle Scholar
Wilson, M (1989) Igneous Petrogenesis. London: Unwin Hyman, 466 pp.CrossRefGoogle Scholar
Wilson, M and Patterson, R (2001) Intra-plate magmatism related to hot fingers in the upper mantle: evidence from the Tertiary-Quaternary volcanic province of western and central Europe. InMantle Plumes: Their Identification through Time (eds R Ernst and K Buchan), pp. 37–58. Geological Society of America Special Paper 352.Google Scholar
Xu, YG, Ma, JL, Frey, FA, Feigenson, MD and Liu, JF (2005) Role of lithosphere–asthenosphere interaction in the genesis of Quaternary alkali and tholeiitic basalts from Datong, western North China Craton. Chemical Geology 224, 247–71.CrossRefGoogle Scholar
Yeganeh, TM, Torkian, A, Christiansen, EH and Sepahi, AA (2018) Petrogenesis of the Darvazeh mafic-intermediate intrusive bodies, Qorveh, Sanandaj-Sirjan zone, Iran. Arabian Journal of Geosciences 11, 1–20.CrossRefGoogle Scholar
Ying, J, Zhang, H, Tang, Y, Su, B and Zhou, X (2013) Diverse crustal components in pyroxenite xenoliths from Junan, Sulu orogenic belt: implications for lithospheric modification invoked by continental subduction: Chemical Geology 356, 181–92.CrossRefGoogle Scholar
Yu, SY, Chen, LM, Lan, JB, He, YS, Chen, Q and Song, XY (2020) Controls of mantle source and condition of melt extraction on generation of the picritic lavas from the Emeishan large igneous province, SW China. Journal of Asian Earth Sciences 203, 104534.CrossRefGoogle Scholar
Zeng, G, Chen, L-H, Xu, X-S, Jiang, S-Y and Hofmann, AW (2010) Carbonated mantle sources for Cenozoic intra-plate alkaline basalts in Shandong, North China. Chemical Geology 273, 35–45.CrossRefGoogle Scholar
Zeng, G, Chen, LH, Hofmann, AW, Jiang, SY and Xu, XS (2011) Crust recycling in the sources of two parallel volcanic chains in Shandong, North China. Earth and Planetary Science Letters 302, 359–68.CrossRefGoogle Scholar
Zhang, SQ, Mahoney, JJ, Mo, XX, Ghazi, AM, Milani, L, Crawford, AJ, Guo, TY and Zhao, ZD (2005) Evidence for a widespread Tethyan upper mantle with Indian-Ocean-Type isotopic characteristics. Journal of Petrology 46, 29–58.CrossRefGoogle Scholar
Zindler, A and Hart, S (1986) Chemical geodynamics: Earth and Planetary Science 14, 493–571.CrossRefGoogle Scholar
Zou, H, Zindler, A, Xu, X and Qi, Q (2000) Major, trace element, and Nd, Sr and Pb isotope studies of Cenozoic basalts in SE China: mantle sources, regional variations, and tectonic significance. Chemical Geology 171, 33–47.CrossRefGoogle Scholar
Figure 0

Fig. 1. Late Cenozoic volcanic centres and active faults in Iran. The study area is included in the red rectangle (the rectangle shows the study area from Allen et al. 2013).

Figure 1

Fig. 2. Schematic geological map showing the location of lava flows in the study area. After Emami et al. (1993). The sampling sites are shown. Abbreviations in the diagram are Gh (Ghezelche-Kand), IL (Illanlu), A (Ahmad-Abad), T (Tahmoures) and G (Ghare-Toreh).

Figure 2

Fig. 3. Overview of the products object of this study. (a) Scoria cones, (b) volcanic bombs, (c) scoria and lavas and (d) gneissic xenoliths.

Figure 3

Fig. 4. Representative photomicrographs of the studied volcanic rocks. (a) Microlithic porphyry texture; (b) glomeroporphyry texture; (c) porphyry texture; (d) sieve texture of clinopyroxene; (e) amphibole; (f) skeletal olivine. Ol: Olivine; Cpx: Clinopyroxene; Qtz: Quartz; Bt: Biotite; Pl: Plagioclase.

Figure 4

Fig. 5. (a) Classification scheme of Morimoto et al. (1988), showing that the pyroxenes are diopside–salite in composition; (b) backscattered electron image microphotograph of clinopyroxene from Illanlu (IL-C7); and (c) mineral compositional variation from rim to rim.

Figure 5

Fig. 6. (a) Variations of forsterite content in olivine vs Mg# of whole rock. Olivines in equilibrium with the host lavas plot between the dashed lines; (b) Na2O + K2O (wt %) vs SiO2 (wt %) diagram for the QBB rocks (Le Bas et al.1986).

Figure 6

Fig. 7. Bivariate diagrams of selected major and trace elements against MgO (wt %).

Figure 7

Fig. 8. (a) Chondrite-normalized REE diagram and (b) primitive-mantle-normalized trace element diagram for QBB rocks. Normalization values from Sun & McDonough (1989).

Figure 8

Table 1. Rb–Sr and Sm–Nd isotopic data for the Qorveh–Bijar basaltic rocks

Figure 9

Fig. 9. (a) 87Sr/86Sr vs 143Nd/144Nd. The compositions of Enriched Mantle 1 (EM1) and Enriched Mantle 2 (EM2) reservoirs come from Zindler & Hart (1986). The fields for OIBs and high-μ (HIMU) come from Hart (1988), and the composition of gneissic xenoliths comes from Azizi et al. (2014). (b) The Mg# of clinopyroxene is plotted against the value of cpx-meltKdFe–Mg (Putirka et al. 2003; Putirka, 2008). Values of cpx-meltKdFe–Mg closely match both the equilibrium ranges of 0.27 ± 0.03 and 0.28 ± 0.08 indicated by Putirka et al. (2003) and Putirka (2008), respectively.

Figure 10

Table 2. Results from clinopyroxene-based thermobarometery, olivine-based thermometry for QBB rocks

Figure 11

Fig. 10. (a) Y (ppm) vs Rb (ppm) and (b) Sm vs Rb diagrams for investigating the fractional crystallization for the QBB rocks; (c) Th/Yb vs 87Sr/86Sri and (d) Nb/Th vs ϵNd for studied rocks.

Figure 12

Fig. 11. Geochemical evidence for crustal contamination in Qorveh–Bijar lavas. (a) AFC modelling for the QBB rocks (gneissic xenoliths of Kord, 2012); (b) 143Nd/144Nd vs Sr (Moghadam et al. 2014); (c) 87Sr/86Sr vs Th (ppm) (Ersoy et al. 2012) with modelled assimilation – fractional crystallization pathways (AFC; r = 0.4 curve); (d) Ba/La vs Ce/Pb and (e) Th/Yb vs Ba/La plots to assess the effects of subducted slab materials on the mantle source of QBBs.

Figure 13

Fig. 12. (a–c) Bulk analyses of QBB mafic rocks display geochemical signatures consistent with melts of a pyroxenite mantle source. (d) Cr (ppm) (a proxy for melting degree) vs normative olivine content of mafic lavas (a proxy for melting pressure) diagram indicates lithospheric drip melting (e.g. Holbig & Grove 2008; Furman et al. 2016). Mafic lava sequences formed through adiabatic upwelling of asthenosphere define a negative correlation. (e) La/Yb vs Sm/Yb melt curves obtained using the non-modal batch melting equation of Shaw (1970) and mantle sources of both garnet-bearing pyroxenite (Cpx60 + Grt40) and spinel-bearing peridotite (Ol80 + Cpx16 + Grt4) (Kelemen et al. 2003; Pertermann et al. 2004). For both models, the values of C0 come from the primitive-mantle data reported in Sun and McDonough (1989). The QBB rocks plot close to the garnet-bearing pyroxenite-melting curve with ∼1 % of partial melting. Tick marks and numbers along the curves show the degree of partial melting for a given mantle source.

Supplementary material: File

Salehi et al. supplementary material

Tables S1-S5

Download Salehi et al. supplementary material(File)
File 79.8 KB